ORIGINAL RESEARCH PAPER

Vertical tectonics of the High Plateau region, Manihiki Plateau,Western Pacific, from seismic stratigraphy

Huirong-Anita Ai Æ Joann M. Stock ÆRobert Clayton Æ Bruce Luyendyk

Received: 2 May 2006 / Accepted: 19 January 2008 / Published online: 27 February 2008

� The Author(s) 2008

Abstract The Manihiki Plateau is an elevated oceanic

volcanic plateau that was formed mostly in Early Creta-

ceous time by hotspot activity. We analyze new seismic

reflection data acquired on cruise KIWI 12 over the High

Plateau region in the southeast of the plateau, to look for

direct evidence of the location of the heat source and the

timing of uplift, subsidence and faulting. These data are

correlated with previous seismic reflection lines from

cruise CATO 3, and with the results at DSDP Site 317 at

the northern edge of the High Plateau. Seven key reflectors

are identified from the seismic reflection profiles and the

resulting isopach maps show local variations in thickness

in the southeastern part of the High Plateau, suggesting a

subsidence (cooling) event in this region during Late

Cretaceous and up to Early Eocene time. We model this as

a hotspot, active and centered on the High Plateau area

during Early Cretaceous time in a near-ridge environment.

The basement and Early Cretaceous volcaniclastic layers

were formed by subaerial and shallow-water eruption due

to the volcanic activity. After that, the plateau experienced

erosion. The cessation of hotspot activity and subsequent

heat loss by Late Cretaceous time caused the plateau to

subside rapidly. The eastern and southern portions of the

High Plateau were rifted away following the cessation of

hot spot activity. As the southeastern portion of the High

Plateau was originally higher and above the calcium car-

bonate compensation depth, it accumulated more sediments

than the surrounding plateau regions. Apparently coeval

with the rapid subsidence of the plateau are normal faults

found at the SE edge of the plateau. Since Early Eocene

time, the plateau subsided to its present depth without

significant deformation.

Keywords Cretaceous � Hotspot � Isopach map �Manihiki Plateau � Seismic reflection

Introduction and tectonic background

The Manihiki Plateau is an area of elevated ocean floor in

the Western Pacific Ocean. The water depth of the plateau

is at approximately 2.5–3 km, several kilometers shallower

than the surrounding oceanic basins (Fig. 1). Winterer

et al. (1974) divided the Manihiki Plateau into three

regions. The High Plateau, in the southeastern section of

the Manihiki Plateau, is the shallowest and flattest, with

water depths of 2.5–3.0 km (Fig. 1). Up to 1 km of pelagic

sedimentary rock covers the igneous basement. The High

Plateau is bordered by linear depressions thought to be

fault troughs (Winterer et al. 1974). The Western Plateaus

comprise smaller ridges and seamounts northwest of the

High Plateau. The North Plateau is a small and nearly

separate shallow plateau at the northern end of the region.

The edges of the High Plateau are steep and linear along

the southwest, east and northwest, where the plateau has

borders with surrounding basins (Fig. 1). The shallowest

H.-A. Ai � J. M. Stock (&) � R. Clayton

Division of Geological and Planetary Sciences, California

Institute of Technology, Pasadena, CA 91125, USA

e-mail: jstock@gps.caltech.edu

H.-A. Ai

e-mail: huirong.ai@gmail.com

R. Clayton

e-mail: clay@gps.caltech.edu

B. Luyendyk

Department of Geological Sciences, University of California,

Santa Barbara, CA, USA

e-mail: luyendyk@geol.ucsb.edu

123

Mar Geophys Res (2008) 29:13–26

DOI 10.1007/s11001-008-9042-0

and most consistently steep edge, on the east, is defined by

the Eastern Scarp, which bounds a linear, relatively flat-

topped ridge trending NNE that is 500–600 m higher than

the adjacent High Plateau. This is the easternmost of a set

of parallel structures that separate the High Plateau from

the Penryhn Basin 3 km below. It is part of the longer

Manihiki Scarp that borders the entire plateau on the

eastern side (Larson et al. 2002). Towards the southwest of

its center, the High Plateau gradually deepens. The south-

western edge of the High Plateau is bounded by the

Suvarov Trough. This feature is a simple asymmetric

graben that separates the High Plateau from a deeper pla-

teau province (the Nassau Plateau). Along this boundary,

there is no adjacent high standing ridge such as that seen

along the Eastern Scarp. There is some evidence of long-

wavelength folds lower on the Nassau Plateau near the

Samoan Basin. It is suggested that the Nassau Plateau is

deformed by gravity tectonics of Late Cretaceous or Early

Cenozoic age, and that the Suvarov Trough is a break-away

structure (Stock et al. 1998).

Originally it was thought that the High Plateau is a

subsided micro-continent (Heezen et al. 1966). It is now

accepted that the Manihiki Plateau is anomalously thick

oceanic crust, based on the results of Deep Sea Drilling

Project (DSDP) Site 317 on the plateau (Winterer et al.

1974; Schlanger et al. 1976). The formation of the

-170

-170

-165

-165

-160

-160

-15 -15

-10 -10

-5 -5

5000

40003000

5000

3000

4000

4500

4500

5000

4500

5000

4000

5500

5000

4000

4000

40004000 3500

4500

4500

5000

5000

HIGH

PLATEAU

WESTERN

PLATEAUS

NORTHPLATEAU

Manihiki

Rakahanga

Suvarov

Danger Is.

Nassau

Penr

hyn

Basi

n

Central Basin of the Pacific

North

Tok

elau

Basin

Samoan Basin

Suvarov Trough

Dan

ger I

slan

ds T

roug

hs

317

160 180

-150

-30

0 0

m

Figu

re1

°

°

°

°°

°

°

°

° ° °

Eas

tern

Sca

rp-160

-20

-10

°

Fig. 1 Bathymetric map of

Manihiki Plateau. Bathymetry is

from the GEBCO digital atlas

published by the British

Oceanographic Data Centre on

behalf of IOC and IHO, 1997.

Much of this bathymetry is the

same as in Winterer et al.

(1974). Inset is topography from

satellite altimetry (Smith and

Sandwell 1997)

14 Mar Geophys Res (2008) 29:13–26

123

Manihiki Plateau was related to Early Cretaceous intense

volcanism and/or ridge jumping in the vicinity of major

transform offsets. Winterer et al. (1974) believed that the

formation of the Manihiki Plateau was tectonically related

to a large jump in the ridge–ridge–ridge (RRR) triple

junction from the Nova-Canton Trough to a location south

of the Manihiki Plateau. Lonsdale (1997) proposed that the

Manihiki Plateau rifted from the conjugate Hikurangi

Plateau, now adjacent to and east of New Zealand, in Early

Cretaceous time. This rift might be the same as that dis-

cussed in Sutherland and Hollis (2001), who suggested a

Cretaceous spreading ridge south of the Pacific–Phoenix

Ridge at the Gondwana margin. They also suggested that

material accreted from the oblique subduction of the Moa

plate is now part of the Eastern province of New Zealand.

Larson et al. (2002) proposed that during 121–119 Ma the

triple junction originated at the northwest corner of the

original Manihiki Plateau, which was then about twice its

present size. The plateau was subsequently rifted into three

fragments, two of which were rafted away (Stock et al.

1998; Larson et al. 2002). Based on new seafloor fabric

data, Taylor (2006) proposed that Ontong Java–Manihiki–

Hikurangi Plateau was once the single largest oceanic

plateau that has since separated into its various parts.

Geochemical and petrologic analysis of the basement

basalts that have been recovered at DSDP Site 317 con-

firms that the Manihiki Plateau has the composition of

oceanic crust. In major element composition, these base-

ment basalts are of MORB-like tholeiitic type (Schlanger

et al. 1976). Mahoney and Spencer (1991) measured the

isotopic ratios of Pb, Nd and Sr for the basalts of the

Manihiki Plateau and found that most of the samples have

relatively low 206Pb/204Pb, high 208Pb/204Pb, and moderate87Sr/86Sr indicative of mantle sources. They concluded that

the sources of the basalts fall well within the modern-day

oceanic island or hotspot field. Beiersdorf et al. (1995)

surveyed volcanic rocks covering a high seamount NE of

DSDP Site 317, at the eastern edge of the Manihiki Plateau,

just SW of Rakahanga Atoll. Their work confirmed the

results of Mahoney and Spencer (1991), indicating a hot-

spot involvement during the formation of the plateau

basement. Additional analyses of these and other Manihiki

Plateau samples by Ingle et al. (2007) show that several

distinct mantle sources were involved.

Ito and Clift (1998) considered the Manihiki Plateau to

be a Large Igneous Province (LIP) and investigated the

sedimentary record of DSDP Site 317 on the plateau to

determine the thermal subsidence history. They concluded

that the thermal subsidence history of the plateau was

comparable to that of normal seafloor, but less than that

predicted by thermal models of hotspot-affected litho-

sphere. They explained this by invoking late stage

underplating and proposed that the formation of the

Manihiki Plateau was related to a hotspot. The more recent

analysis by Clift (2005) argues that the subsidence is less

than typical for normal oceanic lithosphere, similar to that

reconstructed for the Ontong–Java Plateau.

Prior to coring operations at DSDP Site 317, several

single-channel seismic reflection lines were collected over

the plateau during expedition CATO 3 (Winterer et al.

1974). Considerable research has been done on samples

from DSDP Site 317 in terms of paleontology, geochem-

istry, lithology, etc, as well as analysis of the seismic

reflection data of CATO 3 to constrain the evolution his-

tory of the Manihiki Plateau. The present model for the

tectonic evolution of the Manihiki Plateau is based on

either the low resolution single channel seismic reflection

data collected three decades ago (Winterer et al. 1974), or

simply geochemical and paleontological analysis from the

DSDP drilling sites (e.g. Schlanger et al. 1976; Ito and

Clift 1998). However, the coverage of CATO3 lines is

limited and no attempt has been made to correlate the

drilling results with the seismic reflection data across the

whole plateau. Furthermore, reconstruction of a subsidence

history of the Manihiki Plateau based on DSDP Site 317

samples may not represent the evolution of this whole

region since the drilling site is in a basement depression.

Our regional seismic grid allows us to examine the lateral

variation of seismic facies and thus assess possible varia-

tions in subsidence across the region.

In this paper, we augment the data of Winterer et al.

(1974) by adding new single channel seismic reflection

data and correlate these with previously available seismic

data and DSDP drilling results to constrain the tectonic

subsidence history of the Manihiki Plateau. The study

region in this paper is confined to the High Plateau. To

assess the lateral variability of the stratigraphy of the

Manihiki Plateau, we have constructed isopach maps from

the available seismic reflection data and have established a

more detailed subsidence history model, which shows the

High Plateau has subsided more than adjoining regions.

Data and methods

About 4,000 km of single channel seismic reflection,

SeaBeam swath bathymetry and magnetic field data were

acquired in May 1998 over the High Plateau and sur-

rounding ocean regions as part of the KIWI 12 cruise on

the R/V Roger Revelle (Stock et al. 1998). The seismic

reflection survey utilized two 210 cubic inch air guns. For

regular profiling both guns used a firing interval of 10 s.

The recording length of regular profiles was 7 s. The

average trace interval spacing was 45 m.

The new KIWI 12 seismic reflection data are analyzed

and correlated with the 1972 CATO 3 seismic lines and

Mar Geophys Res (2008) 29:13–26 15

123

with the drilling and seismic results from DSDP Site 317

(Schlanger et al. 1976) (Fig. 2). DSDP Site 317 is located

on a CATO 3 seismic line, so its stratigraphy can be cor-

related directly with the reflection profiles. At DSDP Site

317, seven seismic units and two major unconformities U1

(Early/Late Eocene) and U2 (Late Cretaceous/Early

Eocene) were biostratigraphically dated and tied to the

CATO 3 seismic reflection profiles (Schlanger et al. 1976).

The new KIWI 12 tracks are tied to the existing CATO 3

tracks at three of four crossing points in the study area

shown in Fig. 2. The seismic reflection profiles of the

Glomar Challenger, CATO 3, and the new data, KIWI 12,

are correlated with the section drilled at DSDP Site 317

(Fig. 3). Based on the drilling results (Schlanger et al.

1976), three lithologic Mega-Sequences with seven units

were identified, which are bounded in the profiles by seven

reflectors Q-R1 through eC-R7, from seafloor down to the

basement (Table 1). Figure 4 is an example of the reflec-

tors picked from one of the KIWI 12 seismic profiles.

This study depends on how well the new seismic

reflection profiles are correlated with the old seismic

reflection profiles and with DSDP Site 317. The travel

times of the major reflectors correlate well at tie points E, F

and G between the KIWI 12 and CATO 3 lines (Table 2).

At other crossing points on the High Plateau no clear

reflectors are visible, either because of proximity to the

edge of the High Plateau, or else because the resolution of

the earlier seismic data, CATO3, is poor.

The seven stratigraphic reflectors were then used to

construct isopach maps (in Two-Way-Travel-Time) of

layer thickness and depth below sea surface for each

reflector. These are discussed in detail in the next section.

Discussion of isopach mapping

Previous constraints on stratigraphy and lithology

The lithologic units penetrated in DSDP Site 317 include

910 m of sedimentary rocks and 34 m of underlying

basaltic basement (Schlanger et al. 1976). The sedimentary

rocks are divided into three stratigraphic Mega-Sequences

bounded by unconformities. Mega-Sequence 1, which is

above reflector eE-R5 (U1), is upper Eocene through recent

oozes and chalks; Mega-Sequence 2 is lower Eocene cherty

chalk, underlain by a large unsampled section, and

bounded by reflectors eE-R5 (U1) and lC-R6; Mega-

Sequence 3 is largely Maestrichthian (*67 Ma) to Aptian

(*120 Ma) limestone and volcaniclastic sedimentary

rocks, bounded by reflectors lC-R6 and eC-R7 (Table 1,

Fig. 3). U2 lies at or slightly above reflector lC-R6. In

DSDP Hole 317A, the upper part of the volcaniclastic

section lies below the top of the L. cabri zone which is

dated at 120 Ma (Erba 2004). Geochemical, sedimento-

logical, and paleontological analysis of the volcaniclastic

strata and uppermost basement rocks from DSDP Site 317

shows that they were originally deposited in a shallower

water environment than at present depth, about 200 m

(Schlanger et al. 1976; Jenkyns 1976).

The top of the basalt basement is represented by

Reflector eC-R7. 40Ar/39Ar dating of the basalt from DSDP

Site 317 indicated that the basalts were erupted around

123.7 ± 1.5 Ma (Mahoney et al. 1993; Larson and Erba

1999). The vesicular character of the basalts also indicates

that the eruptions took place in shallow water (Schlanger

et al. 1976).

The seven reflectors and isopach maps

The seven reflectors are correlated across the entire High

Plateau region on the KIWI 12 and CATO 3 data. Reflector

eC-R7 is the deepest reflector identifiable and is interpreted

as the top of the basalt basement. It is traceable across most of

the High Plateau. The 0.6 s two-way-travel-time (TWT)

reflector lC-R6 is suggested to represent either sediment in

Core 33-317-4, the change from Maestrichthian chalk

(*67 Ma) to fairly hard Campanian claystone (*75 Ma)

around 70 Ma, or the hiatus between Core 33-317-6 and

Core 33-317-7, around 89 Ma, where a major lithologic

195°

195°

196°

196°

197°

197°

198°

198°

199°

199°

200°

200°

-15° -15°

-14° -14°

-13° -13°

-12° -12°

-11°° -11°

-10° -10°

-9° -9°

-200

-150

-100

-50

0

50

100

150

200

mgal

gravity

Figure 6

Figu

re

7

Figure 8

Figure 3

Figu

re 4

ManihikiPlateau

DSDPSite317

E

F

G

H

M

KIWI12(1998)CATO3(1972)

Figure 5

Fig. 2 Map of High Plateau showing seismic reflection lines

discussed in this paper, and Sandwell and Smith (1997) satellite

gravity field. Blue lines are CATO 3 collected in 1972; red lines are

KIWI 12 collected in 1998; star is DSDP Site 317, located at 11�S,

162�150W; points E, F, G, H are tie points of KIWI 12 and CATO 3

lines. White lines on top of KIWI 12 are locations of seismic profiles

in Figs. 3, 4, 6, 7 and 8. Black points on line of Fig. 4 are where faults

are picked. Black square is location for Fig. 5

16 Mar Geophys Res (2008) 29:13–26

123

change is present (Schlanger et al. 1976). This reflector is at

or slightly below unconformity U2. No previous work has

given a clear interpretation for this unconformity. For con-

venience, reflector lC-R6 and unconformity U2 are grouped

together for the rest of this paper. The lower part of the Mega-

Sequence beneath reflector lC-R6 is traceable across our

whole study area, indicating that the eruption and subsequent

reworking produced widespread volcaniclastic deposits

across the whole High Plateau. The continuity of the U1

unconformity reflector is good but not as strong as U2. It is

associated with an unconformity in the Middle Eocene rocks.

Reflectors lM-R2 through eE-R5 are sub-parallel to the

seafloor, and have only slight regional changes of sub-

bottom time, indicating a relatively stable sedimentary his-

tory since Early Eocene time. These reflectors are weak

compared to reflectors lC-R6 and eC-R7, and sometimes

unidentifiable on the seismic reflection profiles.

Representative Two-Way-Travel-Time contours and

time-isopach maps using the digitized seismic reflection

profiles were made for the region south of 11.5�S, where

there is adequate seismic coverage (Fig. 5). On these

contour and isopach maps we can see many of the

important features of the sedimentary Mega-Sequences.

Reflectors Q-R1 to eE-R5, Mega-Sequence 1 (above U1)

comprising Upper Eocene chalk to Quaternary calcareous

ooze, are very flat in the region bounded by 197� to

198.5�E, 12� to 13�S on the High Plateau (Fig. 5a). The flat

character of reflectors Q-R1 to eE-R5 is interpreted to

result from a relatively stable evolution history of the

Manihiki Plateau since Early Eocene time. The time iso-

pach maps of the layers bounded by reflectors Q-R1

through eE-R5 in the same small region are relatively

constant except towards the edge of the High Plateau

(Fig. 5d). This uniformity shows that the Manihiki Plateau

mainly experienced sedimentation, and possible subsi-

dence, as a whole since Early Eocene time. In contrast, the

Two-Way-Travel-Time contour map of reflector lC-R6

(U2), above the sedimentary and volcaniclastic rocks, has a

different pattern from that of the reflectors Q-R1 to eE-R5

(Fig. 5b). In Fig. 5b, the topography for the same region is

more undulating than it is on maps of reflectors Q-R1

through eE-R5. The contours have two high points; one

Fig. 3 Correlation between DSDP Site 317, CATO 3 seismic

reflection profile and reflectors Q-R1 to eC-R7 on KIWI 12, new

seismic reflection results. Location of KIWI line shown in Fig. 2.

Stratigraphy of DSDP Site 317 and its correlation with CATO 3 is

from Schlanger et al. (1976). Lower case letters ‘‘b,’’ ‘‘c,’’ ‘‘d,’’ and

‘‘e’’ identified by Schlanger et al. (1976) are correlative with 0.07,

0.17, 0.225, and 0.395 s reflectors, respectively; these are reflectors

lM-R2, mM-R3, mM-R4 and eE-R5 according to our definition.

Reflectors we identify (Table 2) have abbreviated labels R1 to R7 in

this figure. Stratigraphic ranges of Mega-Sequences 1, 2 and 3 are

also marked in plot (see text). Left side of plot, from Schlanger et al.

(1976), shows in black the depths from which cores were taken in the

three holes at this site. Uppermost cores 1, 2, and 3 (leftmost column)

are from hole 317; core 1 at 400 m depth, and cores 2–34 (shown in

black in left column) are from hole 317A; cores 1–45 (2nd column,

black continuous region) are from hole 317B

Mar Geophys Res (2008) 29:13–26 17

123

Ta

ble

1D

escr

ipti

on

of

lith

olo

gic

Meg

a-S

equ

ence

sb

ou

nd

edb

yth

ese

ven

refl

ecto

rsa,

and

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elat

ion

wit

hth

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wo

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eat

DS

DP

Sit

e3

17

(in

s)

Lit

ho

log

yM

ega-

Seq

uen

ceB

ou

nd

ing

refl

ecto

rs/

un

con

form

itie

s

Sit

e3

17

(s)

Lo

wer

bo

un

dar

yU

pp

er

bo

un

dar

y

Inte

rnal

con

fig

ura

tio

n

Ag

e

Cal

c.o

oze

Q-R

10

.0–

0.0

7S

ub

ho

rizo

nta

l,

mo

stly

con

form

able

wit

h

seafl

oo

r,

det

erm

ined

by

corr

elat

ing

wit

h

wel

llo

g

Sea

flo

or

Th

intr

ansp

aren

tu

nit

Qu

at.

(1.8

Ma)

Cal

c.o

oze

lM-R

2—

mM

-R3

0.0

7–

0.1

7S

ub

ho

rizo

nta

l,

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stly

con

form

able

wit

hse

aflo

or

Su

bh

ori

zon

tal,

rela

tiv

ely

tran

spar

ent

Mid

-Lat

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ioce

ne

(5.3

–1

0M

a)

Ch

alk

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a-S

equ

ence

1m

M-R

3—

eM-R

40

.17

–0

.22

5S

ub

ho

rizo

nta

lE

arly

-mid

Mio

cen

e

(10

–2

3M

a)

Ch

alk

eM-R

4—

eE-R

5/U

10

.22

5–

0.3

95

An

op

aqu

eu

nit

ov

erly

ing

a

tran

spar

ent

on

e,

mo

der

ate

inte

rnal

lay

ers

Lat

eE

oce

ne–

Ear

ly

Mio

cen

e

(23

–3

4M

a)

Ch

erty

chal

kM

ega-

Seq

uen

ce2

eE-R

5/U

1—

lC-R

6/U

2b

0.3

95

–0

.6M

ost

lyco

nfo

rmab

le

wit

heE

-R5

Co

nco

rdan

to

paq

ue

un

it,

wit

hin

tern

al

refl

ecti

on

s

ov

erly

ing

a

tran

spar

ent

un

it

Ear

lyE

oce

ne

and

old

eru

nsa

mp

led

un

its

(45

Ma

and

old

er)

Lim

esto

ne,

vo

lcan

icla

stic

san

dst

on

e–

silt

sto

ne

Meg

a-S

equ

ence

3lC

-R6

/U2

b–

eC-R

70

.6–

0.8

7D

eep

est

refl

ecto

r

trac

eab

leo

ver

wh

ole

Hig

h

Pla

teau

,cu

tb

y

fau

lts

loca

lly

Str

on

gre

flec

tor

bel

ow

a

tran

spar

ent

un

it,

ov

erly

ing

sev

eral

inte

rnal

refl

ecti

on

s

Th

ick

nes

sh

igh

ly

var

iab

le;

op

aqu

e

un

it;

inte

rnal

lay

ers

cut

by

fau

lts

loca

lly

Mae

stri

chti

anto

Ap

tian

(65

–1

20

Ma)

Bas

alt

eC-R

70

.87

Dee

pes

tre

flec

tor

trac

eab

le,

loca

lly

chao

tic

Vo

lcan

icE

arly

Cre

tace

ou

s

(pre

-12

0M

a)

aR

eflec

tors

are

nam

edfo

llo

win

gth

eco

nv

enti

on

of

May

eret

al.

(19

86)

wh

ere

the

firs

to

ne

or

two

lett

ers

refe

rto

the

age

of

the

refl

ecto

r,e.

g.,

Q,

Qu

ater

nar

y;

eE,

earl

yE

oce

ne;

mM

,m

idd

le

Mio

cen

e,lC

,la

teC

reta

ceo

us;

eC,

earl

yC

reta

ceo

us

bU

2li

esab

ov

ere

flec

tor

lC-R

6,

and

was

form

edaf

ter

refl

ecto

rlC

-R6

18 Mar Geophys Res (2008) 29:13–26

123

near 197.8�E, 12.9�S, and the other to the northwest, at

197.2�E, 12.2�S. Between the two high regions there is a

depression, where the Two-Way-Travel-Time exceeds

3.8 s.

In the same region, the isopach map of Mega-Sequence

2, bounded by reflectors lC-R6 and eE-R5, is different from

that between reflectors Q-R1 and eE-R5 (Fig. 5d vs. e).

Thickness of this unit varies from 0.1 to 0.3 s (Fig. 5e).

This difference in thickness suggests that the Manihiki

Plateau was influenced by tectonic events prior to reflector

lC-R6 time.

The basement surface of the High Plateau (reflector

eC-R7) exhibits significant relief, similar to reflector lC-R6

(Fig. 5c). This has important implications for the paleo-

water depth history. At DSDP Site 317, the paleodepth near

the top of the volcaniclastic section (200 m above the

basement at this site) was about 100 m, as judged from the

molluscan fauna in recovered cores (Kaufmann 1976).

However, DSDP Site 317 was deliberately drilled in a

basement depression, which was evident on the site survey

seismic lines (Schlanger and Winterer 1976). These seis-

mic profiles, as well as our results, suggest that over much

of the High Plateau, and particularly to the SE of DSDP

Site 317, the top of the basement is more than 100 m

shallower than the volcaniclastic strata drilled at DSDP

Site 317. If this basement is the same age or older than the

basement at DSDP Site 317, and if there has not been much

subsequent tilting, it is possible that a considerable portion

of the High Plateau would have been subaerial during or

shortly after its formation. This would provide a likely

location for subaerial erosion to have contributed to the

volcaniclastic sedimentary rocks inferred to be present

Fig. 4 Seismic profile from KIWI 12, showing identified reflectors

Q-R1 to eC-R7. Location of line is shown in Fig. 2. The lC-R6 and

eC-R7 reflectors are affected by normal faults (picked as thick red

lines). Vertical exaggeration is about 20:1. Seafloor relief correlates to

faults due to draping of Mega-sequences 1 and 2 over fault-controlled

relief. Spherical divergence correction and 20–240 Hz band pass filter

have been applied

Table 2 Comparison of Two-Way-Travel-Time (in s) at the tie points (E, F and G) between the CATO 3 and KIWI 12 seismic reflection lines

Reflectors E F G M

CATO3/KIWI12

(161.65�W, 12.65�S)

CATO3/KIWI12

(161.91�W, 13.34�S)

CATO3/KIWI12

(162.39�W, 12.78�S)

CATO3/KIWI12

(162.2�W, 12.1�S)

Q-R1 0.00/0.00 0.00/0.00 0.00/0.00 0.0/0.0

lM-R2 0.05/0.06 0.04/0.06 0.05/0.05 0.07/0.09

mM-R3 0.14/0.15 0.16/0.18 0.12/0.14 0.17/0.18

mM-R4 0.18/0.21 0.23/0.27 0.24/0.20 0.225/0.25

eE-R5 0.33/0.35 0.36/0.36 0.395/0.42

lC-R6 0.39/0.40 0.38/0.38 0.43/0.44 0.6/0.64

eC-R7 n.a./0.54 0.57/0.58 0.68/0.68 0.87/0.87

Mar Geophys Res (2008) 29:13–26 19

123

-163°

-163°

-162°

-162°

-161°

-161°

-13° -13°

-12° -12°

3.6

3.8

4.0

4.2

4.4

4.6

secs

TWTT(s)

4A

A’

-163°

-163°

-162°

-162°

-161°

-161°

-13° -13°

-12° -12°

3.6

3.8

4.0

4.2

4.4

4.6

secs

TWTT(s)

A

A’

-163°

-163°

-162°

-162°

-161°

-161°

-13° -13°

-12° -12°

3.6

3.8

4.0

4.2

4.4

4.6

secs

TWTT(s)

4

4

A

A’

-163°

-163°

-162°

-162°

-161°

-161°

-13°

-12°

0.05

0.10

0.15

0.20

0.25

0.30

0.35

0.40

TWTT(s)

0.4A

A’

-163°

-163°

-162°

-162°

-161°

-161°

-13°

-12°

0.05

0.10

0.15

0.20

0.25

0.30

0.35

0.40

TWTT(s)

0.10.

1

A

A’

-163°

-163°

-162°

-162°

-161°

-161°

-13°

-12°

0.05

0.10

0.15

0.20

0.25

0.30

0.35

0.40

TWTT(s)

0.2

A

A’

(a) Depth of R5

(b) Depth of R6

(c) Depth of R7

(d) R5-seafloor

(e) R6-R5

(f) R7-R6

Fig. 5 (a) Two-Way-Travel-Time (TWTT) contour map showing

time-depth of the eE-R5 reflector below sea surface. Line AA0 is

section line for Fig. 10; (b) Two-Way-Travel-Time contour map of

lC-R6 below sea surface; (c) Two-Way-Travel-Time contour map of

eC-R7, top of basement; (d) Time-isopach map of thickness of layer

bounded by reflectors eE-R5 and Q-R1; (e) Time-isopach map of

thickness of layer bounded by reflectors lC-R6 and eE-R5; (f) Time-

isopach map of thickness of layer bounded by reflectors eC-R7 and

lC-R6

20 Mar Geophys Res (2008) 29:13–26

123

throughout the plateau, above reflector eC-R7 (Winterer

et al. 1974). In addition to volcaniclastic material produced

in subaerial eruptions, some of the volcaniclastic strata

could be epiclastic material derived from subaerial base-

ment highs. This would provide a plausible explanation for

the otherwise puzzling occurrence of a considerable per-

centage of rounded grains in the volcaniclastic sandstones

in DSDP Site 317 (noted by Jenkyns 1976).

Boundaries of the high plateau

Tilted volcaniclastic beds are observed on the western rim

of the High Plateau, near the Suvarov Trough. The upper

Tertiary sedimentary layers are relatively flat (Fig. 6).

Seismic profiles indicate that the sediment deposited in

Suvarov Trough is about 200 m thick (Fig. 6). Whether

these sediments were formed by primary deposition or by

Fig. 6 Western rim near

Suvarov Trough. Locations are

shown in Fig. 2. No spherical

divergence correction, Bandpass

filter same as Fig. 4

Fig. 7 Southwestern edge

showing disturbed High Plateau

relief. Others same as Fig. 6.

See Fig. 2 for location

Mar Geophys Res (2008) 29:13–26 21

123

slumping of the sediments from the High Plateau is not

clear from the profile.

Deeply buried basement highs are clearly visible in seis-

mic reflection profiles from both the CATO 3 and KIWI 12

cruises. The reflectors Q-R1 through eE-R5 are draped over

the previously disturbed High Plateau relief (e.g. Fig. 7).

This basement high identified on the profiles is near the

southwestern edge of the High Plateau (see Fig. 2).

At the eastern edge of the High Plateau, the NNE-

trending Eastern Scarp comprises a zone (100 km wide) of

four to five parallel linear ridges and basins each with a

relief of 500–2,000 m. This ridge and trough region con-

tinues south and north of the High Plateau where it is

named the Manihiki Scarp. Stock et al. (1998) interpret this

feature as a transtensional dextral fracture zone formed

during a change in spreading direction in Late Cretaceous

time. The eastern part of the High Plateau appears trun-

cated with an adjacent block possibly sheared off and

transported away from the High Plateau (Fig. 8). As seen

in Fig. 8, younger sediments are displaced upward, sug-

gesting that the Eastern Scarp was active after these

younger sediments were deposited. On the three crossings

of the Eastern Scarp on the KIWI 12 cruise, as well as on

one crossing obtained on the NBP0304 cruise, multibeam

bathymetric data show that between 12�150S and 13�150Sthis elevated basement rim has a consistent, and relatively

flat top, at a depth of about 2,100 m (Fig. 9). It is possible,

therefore, that this linear ridge was extruded and/or uplifted

when the High Plateau was at a depth of about 700 m. In

this case the linear ridge may have then been eroded to sea

level, after which it would have experienced 2,100 m of

subsidence to bring it down to the depth at which it is found

today.

There are significant normal faults throughout and along

the margins of the High Plateau, cutting seismic reflectors.

For instance, Fig. 4 shows faults displacing reflector

eC-R7, the top of the basement, and the lC-R6 reflector, the

top of the volcaniclastic layer. The faults do not affect the

U2 unconformity. One possible explanation for the origin

of these faults is strain associated with the cooling of the

plateau, after the volcanism stopped. Another explanation

is regional rifting, as explained next.

Joseph et al. (1993) reported faults striking NW–SE on

the northeast and southwest edges of the Manihiki Plateau.

Joseph et al. (1993) related the faults to the mid-Cretaceous

reorganization of the Pacific–Farallon–Phoenix ridge–

ridge–ridge (RRR) triple junction (PAC–FAR–PHN).

Larson et al. (2002) also related the faulting of the Mani-

hiki Plateau to the tectonic evolution of this RRR triple

junction, and proposed that the Manihiki Plateau was

broken and lost two pieces that have drifted away: one

piece to the NE (on the Farallon plate) and one to the

southeast (rifted from the Eastern Scarp onto the Phoenix

plate). An additional piece containing the Hikurangi Pla-

teau rifted away to the south (Lonsdale 1997; Billen and

Stock 2000). Therefore, the widespread normal faults

throughout and along the margins of the High Plateau were

probably the joint result of younger post-volcanic cooling

and the earlier evolution of the triple junction with break-

up of the original Manihiki Plateau.

Very few minor faults cut the upper Cenozoic sedi-

mentary layers, and they are confined to steep escarpments,

Fig. 8 KIWI 12 Seismic

reflection profile records of

Eastern Scarp, High Plateau of

Manihiki Plateau

22 Mar Geophys Res (2008) 29:13–26

123

such as the western edge. These faults may be caused by

slumping near the edge of the plateau and provide little

evidence for the whole tectonic evolution history of the

plateau.

Model for Manihiki Plateau evolution

The schematic cross-section shown in Fig. 10d is derived

from the maps of Fig. 5a–f. As discussed previously, the

Q-R1 through eE-R5 (U1) reflectors are relatively flat,

while the lC-R6 and eC-R7 reflectors show more relief.

Normal faults cut through the eC-R7 and lC-R6 reflectors,

but do not affect unconformity U2 and the section above it.

As a model to explain the formation of the Manihiki

Plateau, we propose that during Early Cretaceous time, the

plateau began to form due to volcanism from a mantle

hotspot, ultimately reaching shallow depths. The presence

of a hotspot is suggested by the isotopic analysis for basalts

of the plateau done by Mahoney and Spencer (1991) and

Beiersdorf et al. (1995), who noted that basalts of the

Manihiki Plateau fall well within the category of oceanic

island or near-ridge hotspots. By 120 Ma, the activity of

the hotspot had created the basement of the original

Manihiki Plateau, and formed the top of the basement. The

uppermost basement basalts at DSDP Site 317 were

erupted in shallow water, while the region of the southeast

High Plateau was shallower and could have been subaerial.

1875 2250 2625 3000 3375 3750 4125 4500 4875 5250 5625Depth (m)

161˚ 30'W

161˚ 30'W

161˚ 15'W

161˚ 15'W

161˚ 00'W

161˚ 00'W

160˚ 45'W

160˚ 45'W

13˚ 30'S 13˚ 30'S

13˚ 15'S 13˚ 15'S

13˚ 00'S 13˚ 00'S

12˚ 45'S 12˚ 45'S

12˚ 30'S 12˚ 30'S

12˚ 15'S 12˚ 15'S

12˚ 00'S 12˚ 00'S

Fig. 9 Bathymetric contour map of SeaBeam data from KIWI 12 and

Simrad EM120 data from NBP0304 showing the relatively flat

character of the high ridge along the Eastern Scarp of the Manihiki

Plateau

Erosion Erosion

EruptionEruption

Triple junction

(a) ~120 Ma

Water

R7

Riftaway

Subside

Collapse

(b) 120~118 Ma

Water

R7

SubsideR7

(c) 118~89(?) Ma

R6

Water

U2

Subside

R7

(d) 89(?) Ma-presentWater

R6

R5

R1(seafloor)

U2

50 km

Fig. 10 Schematic sketch showing evolution of High Plateau region

of Manihiki Plateau based on cross-section AA0 in Fig. 5. Vertical

axis (not to scale) is exaggerated for clarity. (a) During Early

Cretaceous time (*120 Ma), plateau was partly above sea level due

to activity of a hotspot. Basalt basement and volcanic edifices were

formed by eruption. Volcaniclastic sedimentary layers were deposited

by erosion of subaerial basement. Original plateau might have been

near a triple junction; (b) 120–118 Ma: hotspot activity died away

and plateau began to experience cooling and subsidence, as evidenced

by the change in marine fauna. Calcareous sediments were deposited

overlying volcaniclastic layers. During this time, original plateau

might have been separated and rifted away. Normal faults were

formed either from contraction during cooling, or from separation and

rifting of plateau; (c) 118–89(?) Ma: Deposition of calcareous

sediments continued; formation of reflector lC-R6. Faults cut through

eC-R7 and lC-R6 reflectors. Plateau continued subsiding; Unconfor-

mity U2 is probably formed during this time; (d) Since Early Eocene

time, plateau subsided to its present depth

Mar Geophys Res (2008) 29:13–26 23

123

The ancient RRR triple junction would originally have

been northeast of our study area (Larson et al. 2002). The

suggested formation near a RRR triple junction is compa-

rable with the evolution history of Shatsky Rise, located

northwest of the Manihiki Plateau (Ito and Clift 1998).

Volcaniclastic sedimentary layers were deposited on top of

the basement as erosion progressed. A considerable per-

centage of grains were rounded during this erosion and

transportation (Fig. 10a). The deposition of the volcani-

clastic sediments was widespread throughout the plateau,

suggesting that the volcanic source regions were also

widespread.

The short-lived volcanism died away after 118 Ma.

Subsequent lithospheric cooling caused the Manihiki Pla-

teau to subside. Calcareous sediments began to be

deposited at this time (Fig. 10b). The ‘‘shallow-shelf’’

character of the volcaniclastic sedimentary rocks, in con-

trast with the deep water sedimentary rocks overlying the

volcaniclastic rocks, indicates rapid sinking of the plateau

at this time. Several bathymetric depressions were formed

by the rapid subsidence of the plateau. Possibly coeval with

the subsidence are the normal faults that cut through

reflector eC-R7 and the overlying volcaniclastic layers. It

seems likely that these normal faults were, at least in part,

formed during the rapid subsidence of the plateau follow-

ing the end of hotspot activity. The original plateau, which

might have been a very large subaerial volcanic province,

was separated. One section rifted away to the south (the

present Hikurangi Plateau). Another section rifted away to

the east and is now missing. Deposition of limestone and

cherts continued from 118 to *89 Ma when reflector

lC-R6 was formed (Fig. 10c). The undulating geometry of

reflector lC-R6 is the result of either irregular basement

topography, or lateral variation in cooling rates.

Since the location of reflector lC-R6, and/or the

unconformity U2, is not well defined from the present drill

hole data, there is no reliable explanation for its formation

and age. One possibility is given by Schlanger et al.

(1976), who postulated based on analysis of foraminifera in

the sedimentary rocks that an uplift event happened at the

Santonian-Turonian (*89 Ma) boundary to form the

reflector lC-R6. Although uplift of the plateau is consistent

with the paleontological data, it is difficult to explain the

driving force for this uplift. Winterer et al. (1974) sug-

gested that this unconformity was formed after Paleocene

time, but this seems to be ruled out by the results of DSDP

Site 317.

Since Early Eocene time or earlier (starting somewhere

within the sampling gap from 411 to 554 m at DSDP Site

317), thick layers of pelagic sediments were deposited on

the top of the Manihiki Plateau (Fig. 10d). Foraminifera

are very scarce in the clay-rich cherty chalk of the Lower

Eocene, referred to as Mega-Sequence 2 in this work

(Jenkyns 1976). The foraminiferal content and CaCO3

concentrations are known to be indices of carbonate dis-

solution, and the lack of foraminifera suggests a shallow

calcite compensation depth (CCD), which is related to the

productivity and dissolution rate of CaCO3 in the ocean. A

shallow calcite compensation depth is consistent with

results from ODP leg 199 to the east, which showed that

the CCD near the equator was shallower than 3,300 m in

the Paleocene and Eocene, with a sudden deepening to ca.

4,500 m at the Eocene/Oligocene boundary (Rea and Lyle

2005). A deepening of the CCD by ca. 1 km at this time is

found in various other ocean basins, suggesting a global

oceanographic event (Rea and Lyle 2005).

Mega-Sequence 1, the Late Eocene to Recent oozes and

chalks, bounded by reflectors Q-R1 and eE-R5, is charac-

terized by dominantly calcareous strata with carbonate

contents greater than 90%, consistent with deposition

above the CCD. Unconformity U1, between Mega-

Sequences 1 and 2, may be due to submarine erosion

(Jenkyns 1976). It is older than the global deepening of the

CCD that occurred near the Eocene/Oligocene boundary;

however, more drill hole studies in this part of the SW

Pacific are needed to establish its significance. The smooth

pattern of the time-isopach maps of reflectors Q-R1

through eE-R5 indicates little tectonic activity since Late

Eocene time (Fig. 10d).

Discussion

Our results allow us to revisit the conclusions of Ito and

Clift (1998) concerning the history of the Manihiki Plateau.

Ito and Clift (1998) constructed a subsidence curve for the

Manihiki Plateau based on the results of DSDP Site 317.

Our study connects the seismic stratigraphy of this hole to

the region to the southeast, the High Plateau. Although, as

Ito and Clift (1998) summarize, DSDP Site 317 may not

have been subaerial, considerable portions of the plateau to

the southeast may have been near or above the sea surface,

judging from the basement relief visible in the seismic

sections. This provides a likely explanation for the com-

mon occurrence of rounded grains in the sandstones drilled

in the volcaniclastic section at DSDP Site 317.

The plate tectonic model of Larson et al. (2002) and

Taylor (2006) would allow for the possibility that the

plateau was originally much larger. In this scenario, very

large pieces of the plateau rifted off to the east (now sub-

ducted?) and to the south to form the Hikurangi Plateau as

suggested by Lonsdale (1997). In such a case, modeling of

the load due to the hotspot volcanism and the predicted

subsidence is more complicated, because the original load

could have included a much larger mass than what

presently comprises the Manihiki Plateau. This was not

24 Mar Geophys Res (2008) 29:13–26

123

envisioned in the subsidence calculations of Ito and Clift

(1998). Areas that are now rafted away could have pro-

vided a much more extensive source of subaerial

weathering of the volcanic edifices, to produce the thick

volcaniclastic section preserved above the basement in the

High Plateau.

Recent observations of the Hikurangi Plateau adjacent

and east of New Zealand (Hoernle et al. 2004) are also

relevant to the interpretation of the Manihiki Plateau area,

because of the proposed original connection between the

two plateaus prior to rifting along the seafloor spreading

system that formed the Osbourn Trough (Lonsdale 1997;

Billen and Stock 2000). On the basis of swath bathymetry

and other underway geophysical data from the Hikurangi

Plateau, Hoernle et al. (2004) identify several major stages

of volcanism in its history: first, overall plateau growth;

second, late-stage isolated volcanic cones; third, intrusions

along normal faults that accommodated the separation of

the Hikurangi Plateau from the Manihiki Plateau. Their

model requires the surface of the Hikurangi Plateau to

subside from near sea level to about 1,500 m water depth

before the eruption of the late-stage volcanoes, and prior to

the separation of the Hikurangi and Manihiki Plateaus.

Hoernle et al. model would suggest that the Manihiki and

Hikurangi Plateaus would have been adjacent to one

another until after the 1,500 m of subsidence had occurred.

In our study area on the High Plateau of the Manihiki

Plateau, we do not see evidence for late stage seamounts of

the type identified by Hoernle et al. (2004) on the Hiku-

rangi Plateau. However, we only have swath bathymetry

coverage along our track lines, which cover a very limited

area of the High Plateau. We identify basement highs along

some of the steep bathymetric scarps on the Eastern Scarp

and the Suvarov Trough (Fig. 10). There are also some

low-relief highs within the plateau interior, which are now

nearly buried by the Tertiary sedimentary rocks. These

highs consist of unstratified, most likely igneous, material,

and in some cases we cannot rule out the possibility that

these represent a later stage of volcanism than that which

formed the major part of the plateau. The high ridge along

the Eastern Scarp could have formed by magmatism along

a marginal fault system, similar to the suggested origin of

Hoernle et al. (2004) for some of the ridges on the Hiku-

rangi Plateau. Its roughly flat-topped character, if due to

erosion at sea level, would require it to have been emplaced

after the main plateau had already sunk to depths in excess

of 700 m. At the eastern edge of the High Plateau, conical

seamounts such as Mt. Eddie and Rakahanga Atoll were

surveyed and sampled. Petrography and geochemistry

analysis of the basement samples suggest subaerial volca-

nic origin, and isotopic dating shows that age of these

seamounts is in accord with the formation of the Manihiki

Plateau (Beiersdorf et al. 1995).

We do not see a straightforward way to reconcile the

models of Larson et al. (2002) and Hoernle et al. (2004) in

terms of the timing of separation of the three hypothetical

former plateau fragments. Larson et al. (2002) suggest that

the Eastern Scarp rifted at *119 Ma, but Hoernle et al.

model would require somewhat later timing for the sepa-

ration between the Manihiki and Hikurangi Plateaus

because the plateaus would have to subside to *1,500 m

water depth prior to separation. In this case, one might

expect a continuation of the Eastern Scarp to be present on

the east side of the Hikurangi Plateau, but no such scarp is

visible in existing bathymetric data.

There is currently no precise constraint on the time of

separation of the Manihiki and Hikurangi plateaus because

the intervening seafloor formed by the Osbourn Trough

spreading system appears to have been all formed during

the Cretaceous Normal Superchron (120–84 Ma) and lacks

magnetic reversal isochrons. To have the Manihiki–Hiku-

rangi rifting be later than 119 Ma would not violate the

marine geophysical data. However, to invoke diachronous

rifting of the various plateau segments leads to geometrical

complications in the implied triple junction configurations

in the region.

Conclusions

Seven seismic reflectors are traced across the southeastern

part of the High Plateau region of the Manihiki Plateau.

Isopach maps of the time thickness of layers bounded by

pairs of reflectors indicate unequal thinning and thickening

during Early to Late Cretaceous time in the studied area. In

contrast, the thickness of layers bounded by reflectors

formed since Early Eocene time is relatively constant in the

same region. A model to explain this observation is that a

thermally elevated plateau rapidly subsided when the hot-

spot died away during Early Cretaceous time. Differential

subsidence caused the local thickening observed in the

isopach maps and may be related to the extensive set of

normal faults cutting through the basement and the Lower

Cretaceous volcanic sedimentary rocks found at the

southeast edge of this area. The activity of the hotspot was

in a near-ridge environment during the formation of the

original Manihiki Plateau. The Eastern, or Manihiki, Scarp

is related to the rearrangement of the PAC–FAR–PHN

triple junction near the Manihiki Plateau around 121–

119 Ma (Stock et al. 1998; Larson et al. 2002) and marks

where a piece of the High Plateau rifted away. The plateau

began to subside to its present depth after Early Cretaceous

and before Early Eocene time. The relatively constant

pattern of sediment deposition since Early Eocene time

indicates that the plateau experienced little tectonic evo-

lution except thermal subsidence since then.

Mar Geophys Res (2008) 29:13–26 25

123

Acknowledgments We thank Patricia Persaud for her insightful

discussion and numerous suggestions. We also thank graduate and

undergraduate students from Caltech and UC Santa Barbara for their

assistance on KIWI 12. Research supported by University of

California San Diego, Scripps Institution of Oceanography, UC Santa

Barbara, and Caltech. Nathaniel B. Palmer multibeam data collection

was supported by NSF OPP-0126334. Contribution 8945, Division of

Geological and Planetary Sciences, California Institute of Technol-

ogy. Contribution of the Institute for Crustal Studies, University of

California, Santa Barbara, number 0810.

Open Access This article is distributed under the terms of the

Creative Commons Attribution Noncommercial License which per-

mits any noncommercial use, distribution, and reproduction in any

medium, provided the original author(s) and source are credited.

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